Glaciers are composed of four components: ice, water, air, and rock debris. Glacier ice is a unique form of natural ice that is dynamic; it moves internally and over the substratum and is thus a powerful agent for modifying the landscape. Glacier ice is fundamentally different from other kinds of ice, such as sea ice, lake ice, permafrost, etc. Ice is the main component of all glaciers, and ice has several special properties that are important for understanding how glaciers operate.

First, ice behaves as a plastic material under relatively low pressure. This accounts for internal deformation and flowage that takes place in all glaciers. The hardness of ice varies with temperature. At 0°C, ice has a hardness of 1.5 on the Mohs scale; at -70°C the hardness is 6 (Nesje and Dahl 2000). However, temperature below -30°C is rarely achieved in glaciers. Next ice is subject to pressure melting In other words, ice can be melted at temperatures below 0°C by application of pressure--see Fig. 2-1. Because of this and because water has a higher density than ice, water is able to exist under high pressure beneath many glaciers. Finally ice and water have relatively high heat capacities, and phase changes between ice and water involve large heat transfers. Thus the nature of a glacier depends in large part on its thermal regime.

Glacier ice is formed by a process of snow metamorphism, which may be accompanied by seasonal melting and refreezing in some glaciers. The transition from new-fallen snow to glacier ice takes place over a period of years as the snow becomes buried and gradually recrystallized. The process is evident in three main stages:

The metamorphism of snow to ice normally takes about 20-30 years in small alpine or maritime glaciers; however, the process may require a century or more in the interior of large ice sheets. The rate of conversion depends mainly on temperature and amount of annual snow accumulation.

The position on the glacier surface of balance between net accumulation and net ablation is called the equilibrium line altitude (ELA). This line separates the zone of accumulation (névé) above from the zone of melting and calving below. Changes in position of the ELA over time represent changes in the glacier's budget. Increasing height of the ELA indicates a negative budget with ice loss; decrease in ELA represents a positive budget with gain of ice. Nesje and Dahl (2000) have demonstrated that the ELA depends on two climatic factors--winter accumulation and mean summer temperature (at the ELA position). A change in either of these factors will lead to a shift in the ELA.

Several large outlet glaciers descend from the accumulation zone into surrounding valleys. Among these, the Athabasca Glacier and Saskatchewan Glacier have been studied in considerable detail. Both are fed by ice falls, and both exhibit typical (non-surge) flow conditions for alpine valley glaciers--see Fig. 2-6. A remarkable story dramatically illustrates ice movement from the Athabasca Glacier. In April 1965, Julia Oko (age 44) was skiing on the glacier. She apparently broke through an ice bridge and fell 120 feet into a crevasse. She suffered a fractured spine, but was rescued and survived. Eleven years later, her belongings melted out at the glacier snout; these included her skis, rucksack and frozen lunch. At her home, she had the skis and poles mounted in the family room, which she enjoyed until her death in 2005.

The Columbia Ice Field was considerably smaller during the early to mid-Holocene. Wood fragments found at the snout of the Athabasca Glacier are radiocarbon dated at around 8000 years old. The pine (Pinus sp.) and fir (Abies sp.) fragments are thought to represent forest trees that grew on the valley floor now occupied by the glacier (Luckman 1988). Outlet glaciers of the Columbia Ice Field expanded to maximum late Holocene positions circa A.D. 1840-1850, and glaciers remainded extended until the end of the 19th century. During this century, the outlet glaciers of the ice field have retreated significantly.

The Greenland Ice Sheet is scientifically important because it is the only remaining ice sheet in the northern hemisphere, and it is located at the approximate center of ice-sheet development during the Pleistocene. Greenland is relatively accessible, because of the permanent human population located mainly in the southwestern coastal region. The Greenland Ice Sheet represents a bridge to understanding other northern hemisphere ice sheets of the Pleistocene.

A good way to visualize the present state of the ice sheet is to examine the glaciation limit. This represents the lowest elevation at which firn can exist on the ice sheet. This limit is generally lowest in the northwest and northeast (200-400 m) and highest at the southern tip (1600-1800 m). All parts of the ice sheet above the glaciation limit are in the accumulation area (névé). Within the névé, snow exists in various facies depending on its water/ice content--see Fig. 2-7.

Dry-snow facies covers the high, central and northern portions of the ice sheet. Percolation facies covers most of the southern portion and lower northern parts; wet-snow facies is found only in a narrow marginal zone. A narrow marginal zone of the ice sheet is subject to loss, mainly by melting on land and by calving of icebergs where ice reaches the sea in large fjords.

Outlet glacier on Greenland's west coast. This glacier and many others break off
as icebergs, a process called glacier calving. Photograph by Preben Jensen,
Denmark; reproduced here by permission.

View of massive iceberg floating in coastal waters of western Greenland.
Only about 10% of the iceberg's volume is visible, the rest is below sea level.
Photograph by Preben Jensen, Denmark; reproduced here by permission.

The ice sheet occupies a shallow basin that is depressed below sea level and surrounded by mountains--see Fig. 2-8. In effect the ice sheet creates its own topography; it is the highest large plateau in the northern hemisphere. The coldest portion of the ice sheet is the high central zone (below -30°C), with warmer temperatures deeper and toward the margins, where pressure-melting conditions exist at the base of the ice sheet.

Several attempts have been made to estimate the ice-sheet budget. These attempts were difficult in the past because of the great size of the ice sheet and limited information concerning accumulation, melting, and calving of icebergs. Remote sensing observations have improved our ability to gauge changes in the ice sheet. Altimeter measurements from the ERS-1 and ERS-2 satellites documented an overall increase in elevation of the ice surface during the period 1992-2003 (Johannessen et al. 2005). Throughout the vast interior of the ice sheet, surface elevations increased at an average rate of ~6½ cm/year. The narrow marginal zone suffered a decrease at ~2 cm/year. Over the whole of the ice sheet this represents an average increase of ~5½ cm/year, or 60 cm total for the 11-year period of observation. When corrected for crustal depression, the total increase is estimated to be 54 cm. In other words, the Greenland Ice Sheet has gained more than half a meter of surface accumulation overall since the early 1990s, which suggests a strongly positive budget.

Ice movement is generally downward and outward from two major domes in the east-central and southern portions. Movement within most of the accumulation zone is slow, only a few m/year, with movement increasing to about 80 m/year near the firn line. The two domes are separated by a lower saddle from which several fleuves de glâce flow westward. These concentrated, high-velocity ice streams follow subglacial valleys leading westward to Disko and Umanak Bays, where a large volume of icebergs is discharged. Near the calving margin, ice-flow velocities are 2-10 km/year, and much melt water emerges from beneath the ice.

The Greenland Ice Sheet was considerably larger during times of maximum Pleistocene glaciation. It averaged about 500 m thicker and extended to near the edge of the continental shelf. In spite of its greater thickness, the ice sheet did not cover all of the marginal mountains; many nunataks remained in which some plant species may have locally survived glaciation.

On the continental shelf, sea level was some 100-150 m lower, so some banks were exposed as dry land at the edge of the ice sheet. The ice sheet may have been confluent with the Ellesmere Ice Cap to the northwest, although this is a debatable interpretation. During maximum glaciation, the coastal areas were depressed as a result of ice loading on the crust; these areas have now rebounded by 50-200 m.

The ancestral Jakobshavn ice stream expanded several 100 km to the shelf edge on the western side of Greenland at the last glacial maximum ~20,000 years ago (Cofaigh et al. 2013). It began to retreat by 14,800 calibrated-years ago; however, it readvanced significantly during the short-lived Younger Dryas phase about 12,000 calibrated-years ago.

It is generally thought that the Greenland Ice Sheet formed in the late Pliocene or early Pleistocene by coalescence of ice caps and glaciers. New evidence from offshore marine deposits suggests that Greenland had a partial ice cover as far back as seven million years ago (late Miocene). The marine deposits contain large stones transported by drifting icebergs derived from Greenland glaciers.

Antarctica--a mosaic of 40 Galileo images taken through red, green, and violet filters.
Antarctica is about 4000 km across from top to bottom of view. The South Pole is just right
of center, and Ross Ice Shelf is just above the center. Dark zones are open-water seas around
the margin of the ice sheet. Derived from NASA Goddard Space Flight Center, Earth-Galileo
imagery.

Most of the East Antarctic Ice Sheet rests on crust that is above sea level. Recent geodetic surveys have shown that ice at the South Pole moves nearly 10 meters per year toward the northwest (Mullins 1999). Much of the base of the West Antarctic Ice Sheet lies far below present sea level. In other words, the East Antarctic Ice Sheet is land based, whereas the West Antarctic Ice Sheet is marine based. Floating ice shelves and ice tongues have distinctive ice-flow dynamics that involve interaction between the inland ice and the sea.

Marine-based ice sheets are inherently unstable. Such an ice sheet is held in check by back stress from the adjacent ice shelf by shear stress at basal pinning points and along lateral margins--see Fig. 2-10. A slight rise in relative sea level (crustal subsidence), decrease in ice flow, or thinning of the ice shelf would reduce the back stress and could lead to rapid collapse of the ice sheet (Alley 1989). Antarctic ice shelves are sources of enormous tabular icebergs.

The West Antarctic Ice Sheet is drained by a series of ice streams that feed into the Ross Ice Shelf--see Fig. 2-11. These ice streams flow through a region of relatively thin ice (about 1000 m), but they maintain flow rates of around 500 m per year, and they accelerate downstream. They are characterized by highly crevassed surfaces in contrast to the adjacent smooth surface of slow-moving ice. The individual ice streams appear to have complex and different histories of movement. For example, ice stream C seems to have stopped moving around 100 years ago. Most of the rapid motion of ice streams is accomplished by basal sliding over water-lubricated beds or by deformation within a thin, water-saturated sediment layer below the ice (Alley 1989).

The Ross Ice Shelf has fluctuated considerably during the past 45,000 years--see Fig. 2-12. Between 45,000 and 27,000 years ago, the ice edge was quite similar to now, then expanded reaching a maximum position ~13,000 years ago. During the Holocene, the ice margin retreated considerably reaching a minimum position in the middle Holocene between 4000 and 2000 years ago.

Antarctica was glaciated much earlier than was Greenland. A semipermanent ice sheet was established by the end of the Eocene or earliest Oligocene, about 34 million years ago (Ivany et al. 2006). Glaciation in other high-latitude southern lands also started early. Tasmania supported local glaciers in the early Oligocene, around 36 million years ago, when it was a mountainous peninsula of Australia at about 55-56°S latitude (Macphail et al. 1993). The Antarctic ice sheet experienced a major expansion during the early Pliocene, 5-4 million years ago (Bart 2001). This expansion took place in spite of relatively high sea level and warm temperature globally, and the ice volume may have been significantly larger than present ice volume.

The Greenland Ice Sheet has yielded a rich collection of cosmic dust (particles <1 mm in size). About 85% of the cosmic dust particles are carbonaceous chondrites, thought to be samples of the primitive solar system. Cosmic dust falls constantly on the ice sheet and is carried by ice flow to the margin. In some places, ice melts in local basins and the particles settle to the bottom of "blue ice lakes," where they accumulate year after year (Episodes 9, 1986).